The current understanding of the evolution of the atmosphere, hydrosphere, and biosphere on early Earth has been strongly influenced by the following six major paradigms for the geochemical cycles of oxygen, iron, and sulfur: (1) a dramatic change from a reducing to an oxidizing atmosphere at ca. 2.4–2.2 Ga, termed the “Great Oxidation Event” (GOE); (2) Fe-rich oceans until ca. 1.85 Ga; (3) a hydrothermal origin for the global oceanic Fe; (4) SO₄ ²⁻-poor oceans before the GOE; (5) an atmospheric origin for the oceanic sulfur species; and (6) the existence of sulfidic Proterozoic oceans.

Each of the six paradigms has been built on other paradigms, such as those concerning: (1) the behavior of Fe during soil formation, (2) the environments and processes required for the formation of Fe^III oxides in banded iron formations (BIFs), and (3) the origins of siderite and pyrite, as well as (4) the origin of anomalous isotope fractionation of sulfur (AIF-S) in Archean sedimentary rocks. Here, we show that some of the paradigms contradict each other, and that each has serious flaws (contradictions, problems) when they are compared to a variety of observations (geological, mineralogical, or geochemical data from natural samples; laboratory experimental data; results of theoretical studies). In contrast, all of these observations are better explained by the Dimroth-Ohmoto model for Earth's evolution, which postulates that a fully oxygenated atmosphere-ocean system developed by ca. 3.5 Ga.

Examination of the available data from natural and experimental systems has also led us to suggest the following: (1) The geochemical cycles of O, Fe, and S (and other redox-sensitive elements) through the atmosphere–ocean–oceanic crust–mantle–continental crust have been basically the same as today since at least ca. 3.5 Ga. (2) The anaerobic and aerobic microbial biospheres, both in the oceans and on land, developed by ca. 3.5 Ga, playing an important role in the geochemical cycles of nutrients and other elements. (3) The geochemistry of sedimentary rocks (shales, carbonates, cherts) has been basically the same since ca. 3.5 Ga. (4) Fe^III oxides in BIFs were formed by reactions between locally discharged Fe²⁺- and silica-rich submarine hydrothermal fluids and O₂-rich deep seawater. (5) Magnetite in BIFs was formed during high-temperature diagenetic stages of BIFs through reactions between primary goethite or hematite and Fe²⁺-rich hydrothermal fluids. (6) BIFs were formed throughout geologic history. (7) Sulfidic oceans (i.e., the “Canfield ocean”) did not exist during the Proterozoic Eon. However, regional sulfidic seas, like the Black Sea, have existed in globally oxygenated oceans throughout geologic history. (8) The primary carbonate in Archean oceans, as in younger oceans, was Fe-poor calcite.

Furthermore, (9) the pre–1.8 Ga atmosphere was CO₂ rich with the pCO₂ level greater than ~100 present atmospheric level (PAL). CO₂ alone provided the green-house effect necessary to compensate for the young Sun's lower luminosity. (10) The Archean pH values were 4.0–4.5 for rainwater, between 4.5 and 6.0 for river water, and 7.0 ± 0.5 for ocean water. The oceans were saturated with calcite but under-saturated with siderite. (11) The δ¹⁸O of Archean oceans was ~0‰, as today. (12) Fe-rich carbonates (siderite, ankerite) have formed during the diagenesis of sediments throughout geologic history by reactions between the primary calcite and Fe²⁺-rich solutions, either hydrothermal solutions or those derived from biological or abiological dissolution of Fe^III-(hydr)oxides within the sediments.

Other suggestions include: (13) The ranges of δ³⁴S values of pyrite and sulfates in Archean sedimentary rocks are much larger than those quoted in the literature and comparable to those in Proterozoic sedimentary rocks. (14) Pyrites in organic C–rich black shales associated with BIFs were formed by a reaction between Fe²⁺- and SO₄ ²⁻-rich hydrothermal solutions and organic C–rich shales during early diagenetic stages of the host sediments. This reaction also created AIF-S in the pyrite and the residual SO₄ ²⁻. (15) The AIF-S signatures in Archean and younger rocks were not created by the ultraviolet photolysis of volcanic SO₂ in a reducing atmosphere. AIF-S signatures are not evidence for a reducing atmosphere. (16) Contrary to a popular belief that AIF-S–forming events ceased at ca. 2.45 Ga, AIF-S was also formed at later geologic times. (17) The presence of AIF-S signatures in some pre–2.4 Ga rocks, but the lower abundance of AIF-S in post–2.4 Ga rocks, may reflect changes in the mantle-crust dynamics, including changes in the thickness and movements of oceanic lithosphere.